ARTICLE IN PRESS · resulting from plate coupling: (1) the relationship between surface deformation...

Tectonophysics xxx (2009) xxx–xxx

TECTO-124586; No of Pages 18

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Tectonophysics

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Neogene to Quaternary tectonics of the coastal Cordillera, northern Chile

Richard W. Allmendinger a,⁎, Gabriel González b

a Department of Earth and Atmospheric Sciences, Cornell University, Ithaca, NY, USAb Departamento de Ciencias Geológicas, Universidad Católica del Norte, Antofagasta, Chile

⁎ Corresponding author.E-mail address: [email protected] (R.W. Allmending

0040-1951/$ – see front matter © 2009 Elsevier B.V. Adoi:10.1016/j.tecto.2009.04.019

Please cite this article as: Allmendinger, RTectonophysics (2009), doi:10.1016/j.tecto.

a b s t r a c t
a r t i c l e i n f o
Article history:Received 10 July 2008Received in revised form 6 February 2009Accepted 19 April 2009Available online xxxx

Keywords:SubductionTectonicsStructural geologyChileCoseismicInterseismic

The extreme aridity of the northern Chilean Coastal Cordillera enables the complete preservation ofpermanent deformation related to the coupling between Nazca and South America. The region betweenAntofagasta and Arica is characterized by four types of Late Cenozoic structures: EW-striking reverse faults;~NS-striking normal faults; sparse WNWand NNW-striking right-lateral strike–slip faults; and unimodal andbimodal populations of surface cracks. The EW reverse faults occur only between 19° and 21°40' S latitudeand have been active since at least 6 Ma to present. A March 2007 earthquake demonstrates that thisdeformation can occur during the interseismic phase of the plate boundary seismic cycle. The NS-normalfaults are most abundant near Antofagasta and Mejillones and diminish in significance northward. Several ofthese faults, in both the Antofagasta and the Salar Grande area, have been reactivated as reverse faultscausing minor topographic inversion. We suggest that normal faults move during both interseismic andcoseismic deformation and are fundamentally very weak. Surface cracks are ubiquitous throughout theregion and form by both non-tectonic and tectonic mechanisms. A significant proportion forms duringcoseismic deformation and thus may be used to identify long-lived rupture segments on the plate boundary.Overall, forearc deformation is very slow, with permanent strain rates of 1–5 nstrain/year and fault slip ratesless than 0.5 mm/year.

© 2009 Elsevier B.V. All rights reserved.

1. Introduction

The forearc of the Chilean Andes provides an unparalleledopportunity to study the processes that affect the leading edge of amodern continent–ocean convergent plate boundary. This plateboundary periodically generates massive subduction zone earth-quakes which not only affect local communities but also producetsunamis that impact the entire Pacific Ocean rim. The rupture zonefor large and great earthquakes lies beneath the coastal areas innorthern Chile and southern Perú (Dorbath et al., 1990; Tichelaar andRuff, 1991). The earthquakes are the culminations of cycles in whichelastic strains build up over a 100–150 year period as convergencecontinues but the plate boundary is locked (Comte and Pardo, 1991).During the interseismic part of the cycle the Coastal Cordillera isuplifted while the peninsulas of the littoral zone subside; this patternis reversed during the coseismic part of the cycle (Klotz et al., 1999;Pritchard et al., 2002; Chlieh et al., 2004; Pritchard et al., 2006).Deformation features in northern Chile are spectacularly exposedalong the Coastal Cordillera which occupies a position above theseismogenic zone of the Central Andes (Figs. 1, 2).

Linear topographic scarps mark the traces of young faults and foldsthat overlie, and are spatially correlated with, the zone of coupling

er).

ll rights reserved.

.W., González, G., Neogene2009.04.019

between Nazca and South America plates. A major question regardingthese faults, however, is whether the structures move during theinterseismic or coseismic/post-seismic part of the plate boundarycycle. Additionally, surface cracks of both tectonic and non-tectonicorigins are ubiquitous. For the subset of cracks that are tectonic, wecan, likewise, pose the question: are they related to interseismic orcoseismic plate boundary deformation? Long term and large scaleuplift of the coastal line is marked by flights of marine terraces whichare particularly obvious on the Mejillones Peninsula, as well as thecoastal escarpment, itself.

One of the great advantages of the forearc in northern Chile is thehyperarid climate: there is no vegetation to obscure subtle surfacefeatures and no running water to wear them away. The exquisitepreservation of these surface deformation features is virtually uniquein the pantheon of modern forearcs. It is likely that the structuresdescribed here are present to varying degrees in all tectonicallyeroding forearcs; only here, however, are they preserved for study.

In this paper, we focus on the Coastal Cordillera between 24°S and18°S latitude (Fig. 1), a region roughly between 800 and 2000 melevation, bounded to the west by the coastal escarpment and, locally,a narrowmarine terrace, and to the east by the slightly lower elevationlongitudinal valley. We explore two discrete geological aspectsresulting from plate coupling: (1) the relationship between surfacedeformation features and the subduction seismic cycle and (2) theinfluence in the upper crustal deformation regime of the map viewcurvature of the plate coupling zone.

to Quaternary tectonics of the coastal Cordillera, northern Chile,

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2. Tectonic setting

Nearly three decades of intense geophysical investigation hasmade western South America one of the best known convergentmargins on Earth (Bevis et al., 2001; Buske et al., 2002; Götze et al.,1994; Husen et al., 1999; Kendrick et al., 2001; Khazaradze and Klotz,2003; Klotz et al., 1999; Oncken et al., 1999; Pritchard et al., 2002;Sobiesiak et al., 2007; Sobiesiak, 2000; von Huene and Ranero, 2003;Wigger et al., 1994). The current cycle of ocean–continent conver-gence began in the Jurassic following the break-up of the Gondwanasupercontinent (Mpodozis and Ramos,1990) and has been continuingever since with varying degrees of obliquity (Pardo-Casas and Molnar,1987). The remnants of the Jurassic volcanic arc currently composemuch of the basement of the Coastal Cordillera. This ancient arc liestoo close to the modern subduction system (Rutland, 1971); in theAntofagasta area, it is just 75 km east of the trench and 35 km above

Fig. 1. Shaded relief map of northern Chile and adjacent areas; international boundaries shoWadati–Benioff zone from Cahill and Isacks (1992), modified in the zone of interplate couplOncken et al., 1999). The heavy dashed line is the axis of Andean topographic and WBZ sym

Please cite this article as: Allmendinger, R.W., González, G., NeogeneTectonophysics (2009), doi:10.1016/j.tecto.2009.04.019

the subducted plate. This proximity requires that about 200 km offorearc have been removed since the start of the Andean convergentmargin by subduction erosion, an observation that has beenconfirmed by many subsequent studies, which demonstrate the lackof any accretionary prism and document that normal faults arepresent almost all the way to the inner trench wall (e.g., von Hueneand Scholl, 1991; von Huene and Ranero, 2003).

Following a relative lull in convergence rate in the Oligocene andearly Miocene, probably caused by very oblique subduction, orthogo-nal convergence peaked in the early–mid Miocene and has beendecreasing ever since (Pardo-Casas and Molnar, 1987; Somoza, 1998).The current, GPS determined rate of ~6.5 cm/year (Angermann et al.,1999; Kendrick et al., 2003) is less than half of the Miocene rate of~15 cm/year (Pardo-Casas and Molnar, 1987; Somoza, 1998).

During the last two decades, the position of the seismogenic zonein northern Chile has been constrained by using teleseismic data

wnwith long-short dashed line. The long curving lines are contours on the depth to theing by local network and seismic reflection data (Buske et al., 2002; Husen et al., 1999;metry (Gephart, 1994). The labeled boxes show areas described in the text.



Fig. 2. Generalized topographic cross-section showing the relationship of the Coastal Cordillera to the zone of seismic interplate coupling. Modified from Buske et al. (2002).

3R.W. Allmendinger, G. González / Tectonophysics xxx (2009) xxx–xxx

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(Tichelaar and Ruff, 1991) and local networks (Comte et al., 1994;Delouis et al., 1996, 1997; Husen et al., 1999). Recent large thrustearthquakes, including the 1995 Mw 8.1 Antofagasta and the 2007Mw 7.7 Tocopilla events and their after shocks, have shown that theseismogenic zone extends down-dip to a depth of 50±5 km belowthe Coastal Cordillera (Husen et al., 1999, 2000; Buske et al., 2002).The seismogenic zone dips approximately 12°–14° to the east (Fig. 2).The Iquique rupture segment, which extends from Ilo, Peru toMejillones, Chile, has not had a great earthquake since 1877. Modelingof regional campaign GPS data shows that the plate boundary islargely or completely locked during the interseismic part of theearthquake cycle (Norabuena et al., 1998; Bevis et al., 2001; Klotz et al.,2001). InSAR and GPS data combined with elastic dislocation modelsuggest a fully locked zone up to 35 km and a transition zone between35 and 55 km depth (Chlieh et al., 2004); the depth of lockingprobably varies along strike (Khazaradze and Klotz, 2003).

3. Initiation of modern deformation cycle

Fault line scarps disrupt a smooth and, at one time, morecontinuous topography. This modern deformation regime beganafter a long period erosion that culminated in a regional peneplainsurface (the Tarapacá peneplain of Mortimer and Saric, 1972).Recently, surface dating based on cosmogenic nuclides (Carrizoet al., 2008; Dunai et al., 2005) confirms that this peneplain surfacewas probably formed approximately 22–25 Ma before present. Theconsolidation and preservation of the peneplain surface is relativelyclose in time to the onset of the hyperaridity in the Central Andesforearc (Rech et al., 2006). Remnants of the Tarapacá surface can befound in the highest parts of coastal escarpment suggesting that theCoastal Cordillera once extended some distance westward from thepresent day coastline. A U–Th/He transect up the coastal escarpmentnear Tocopilla demonstrates cooling through 40–70 °C at 45±5 Ma(Juez-Larre et al., 2007).

The more recent evolution of the Coastal Cordillera has beencharacterized by coastal escarpment retreat and surface uplift,


terminating in the present day endorheic regime of the westernslope of the Andes, as well as the formation of the Central Depression.Long term hyperaridity may produce intensified plate coupling byreducing the amount of water saturated sediment in the trench (Lamband Davis, 2003). This history begs the question of whether the upliftof the Coastal Cordillera is a consequence of increased plate couplingand therefore if the present day position of the seismogenic zone hasbeen relatively stable since Oligocene–Miocene time.

4. Regional description

In this section, we describe the Late Cenozoic structures of thenorthern Chilean Coastal Cordillera. For the two southern segments,Paposo and Mejillones, we update the observations made by Delouiset al. (1998) with observations of our own; for the rest of the area, weupdate our own work (Allmendinger et al., 2005; Carrizo et al., 2008;Dunai et al., 2005; González et al., 2003a,b, 2006, 2008; González andCarrizo, 2003; Loveless et al., 2005). The Late Cenozoic deformation isoverprinted on a Mesozoic and Paleozoic basement. The mostcommon rock types include highly altered Jurassic volcanics, as wellas red clastic deposits and limestones of Cretaceous age that appear, atleast locally, to have been deposited in extensional half-graben(Allmendinger et al., 2005; Anonymous, 2003). There are alsonumerous Mesozoic intrusive igneous rocks and, locally, Paleozoicmetasedimentary rocks also occur.

4.1. Paposo segment

The Paposo segment (Figs. 1, 3) is one of the topographicallyhighest parts of the Coastal Cordillera. Within this segment, severalfault zones have clear late Cenozoic activity: the Paposo segment ofthe Atacama fault zone and the much less significant Barazarte andSalar de Navidad faults (Arabasz, 1971; Delouis et al., 1998). The lateCenozoic Paposo fault zone is exposed near its southern limit ofexposure where it dips steeply (~75°) eastward and juxtaposes basinfill to the east against basement rocks, or older Cenozoic basin fill to



Fig. 3.Map of the Paposo segments, showing some of the faults and localities describedin the text. Shaded relief and topographic contours the same as in Fig. 1. Dashed whiteline is the 50 km contour of the depth to the subducted plate. Stars show the location ofsubsequent figures.


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thewest, indicating normal faulting (Fig. 4A). The topographic scarp isvery smooth suggesting no significant recent movement in this area.Farther north, incision is insufficient to expose the fault plane but aclear topographic scarp can be seen on the hillside 50–100 m abovethe valley bottom (Fig. 4B). The scarp, which cuts the heads of thealluvial cones, forms a topographic bench on the side of the hill; thedetailed morphology of the cone heads indicates that the valley hasbeen uplifted relative to the mountains, by 2±1 m along a reversefault.

The right lateral, N 20–30° W striking, Barazarte fault has alreadybeen well described by Delouis et al. (1998). New high resolutionimagery permits a more accurate determination of displacement.Headless gullies that truncate against the fault document theprogressive abandonment of channels on the downstream side ofthe main gully as displacement accrued (Fig. 5). Total minimum rightlateral displacement is about 250 m, with 60–100 m of fault offsetbefore the abandonment of each gully. Ridge crests are also offsetacross a subsidiary fault trace. Unfortunately, the landforms arecurrently undated.


The Salar de Navidad fault (Fig. 3) strikes 290°–305° and has alsobeen described by Delouis et al. (1998) and mapped by González andNiemeyer (2005). The dip is not well determined but fractures alongthe main fault trace where it crosses a large valley suggest that itprobably dips 45–60°NE. The 6 m scarp is uplifted on the NE side,indicating a component of reverse movement. The fractures along thefault trace cut the youngest drainages, so there is evidence of relativelyrecent motion. The most recent, and primary, motion on the fault isright-lateral based on consistently offset drainages, small pressureridges along minor left stepping segments of the fault, and theorientation of open fractures in the hangingwall, with strikes between339° and 355°.

4.2. Mejillones segment

The normal faults of the Mejillones segment (Fig. 6) have alreadybeen well described by numerous authors (Armijo and Thiele, 1990;Niemeyer et al., 1996; González et al., 2003a,b; González and Carrizo,2003; González et al., 2006). Here, we emphasize new observationsand analyses that have not received wide distribution.

4.2.1. Faults of the Mejillones PeninsulaNormal faults of the Mejillones Peninsula strike generally NS and

dip to the east, except for the minor Bandurias fault (Fig. 6), whichdips to the west. In general the faults have a listric geometry withgentle rollover anticlines in the hanging walls and unrotatedfootwalls. The rollover geometry on the south side of the peninsulais clearly marked by a gently westward tilted Pleistocene terrace.

The Caleta Herradura normal fault and half graben is the bestexposed of any in the coastal region of northern Chile (Fig. 7). 60–70high sea cliffs exposed the Lower Miocene to Pleistocene strata(Ibaraki, 1990; Koizumi, 1990) which fill the half graben anddemonstrate a protracted history of extension. The westward tiltingand thickening of these strata, combined with the lack of rotation ofthe well preserved Pliocene and younger wave cut terraces and stratain the footwall require that the fault has a listric geometry andshallows with depth. Shear oblique to bedding in the hanging wall(Dula, 1991; Rowan and Kligfield, 1989) is recorded by normal faultsthat dip both east and west in similar numbers and with similar slipmagnitudes (Fig. 8A). We model kinematics of hanging walldeformation with both vertical and synthetic simple shear. With thisinterpretation, the Caleta Herradura listric normal fault shouldbecome horizontal at less than 2 km depth (Fig. 8B). Given about500 m of slip on the Caleta Herradura fault and the fact that thegrowth strata are lower Miocene and younger, we calculate a longterm slip rate of ~0.025 mm/year.

Less is known about the geometry of the Mejillones half grabenbecause of almost non-existent vertical exposure of the hanging wall.The fault also appears to be listric, but it probably shallows tohorizontal at a deeper crustal level given that the half graben is muchwider than the Caleta Herradura half graben. The fault plane, exposedin at least two places, dips ~63° to the east and has down-dipslickensides. It is likewise imaged on high resolution bathymetry andshallow reflection seismic data (Vargas et al., 2005). Alluvial fans fromthe Morro Mejillones form growth strata in the hanging-wall of thisfault, indicating that the most recent continental deposition has beenaccommodated by vertical slip on the Mejillones Fault. Inactivealluvial surfaces, dated at 47,000 years, have been offset across thefault by ~13 m (Marquardt, 2005), yielding a short term slip rate of~0.3 mm/year and a horizontal extension rate of ~0.11 mm/year.

4.2.2. Coastal Cordillera domino blocks, including the Salar del Carmensegment of the Atacama Fault Zone

East of the Mejillones Peninsula, several well known (Niemeyeret al., 1996) and relatively well exposed normal faults dip 60–68° tothe east. These faults strike more northeasterly than many structures



Fig. 4. (A) Outcrop of the Paposo normal fault in late Cenozoic basin fill. (B) Topographic bench produced by reverse fault reactivation of the Paposo normal fault.


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in the area. The conventional interpretation is that the faults aresimply reactivating the Mesozoic Atacama fault zone, but someauthors attribute this change to a change in depth of interseismiccoupling from 35 km to the south to 50 km farther north (Khazaradzeand Klotz, 2003; Loveless, 2007).

Unlike the faults of the Peninsula, both the hanging wall and thefootwall blocks are tilted 4–6° to the west, suggesting a dominonormal fault model (Fig. 9). Using the relationship between tilt andfault dip for horizontal extension from domino blocks (Wernicke andBurchfiel, 1982), we calculate about 4% horizontal extension–or about1200m across the current horizontal distance of 30 km–for this part ofthe Coastal Cordillera. The deposits cut by these normal faults arepoorly dated in the region of study (Naranjo and Paskoff, 1985;


Niemeyer et al., 1996); if we assume an age of initiation of extensionbetween early Pliocene and early Miocene, the resulting horizontalrate of extension is 0.18–0.06 mm/year.

The easternmost fault in this set of domino blocks is the Salar delCarmen fault (Armijo and Thiele, 1990; Delouis et al., 1998; Gonzálezet al., 2003a,b, 2006). The northernmost end of the fault (Fig. 6)exhibits a morphology similar to that seen along the Paposo fault inwhich the scarp forms a subtle bench on the hill slope about 20–40 mabove the topographic break in slope at the base of the hill. This benchhas accumulated a number of small dark clasts that have moved downslope and thus forms a dark band across the hill side (Fig. 10A).Outcrops in a number of the small, steep gullies that cut the scarpshow that the fault dips SE at 65–70° and uplifts a sliver of basement



Fig. 5. Google Earth image of the northern part of the Barazarte fault showing dextral offset of morphological features. (A) Older displaced gullies, (B) and (C) shutter ridges.


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on the southeast with respect to the topographically higher footwall.The major scarp at this locality appears to be inactive: unlike the Salardel Carmen fault farther south, there are no fresh fractures, raw scarpfaces etc. This site, together with the Paposo segment, demonstratesthat the east-dipping fault zones have experienced reverse reactivation(Fig. 10B). The age of the reverse fault motion is younger than the offsetvolcanic ash; elsewhere in the region, similar volcanic ashes havePliocene ages (3.5 and 5 Ma, González et al., 2003a,b).

4.2.3. Open tectonic cracksOpen surface cracks characterize much of the area and are

commonly, but not exclusively, associated with fault scarps as atSalar de Navidad. The ages of these cracks are not known and severallines of evidence reviewed below suggest that the cracks arereactivated multiple times. On the west side of Salar del Carmen,surface cracks were observed to have formed during the 1995 Mw 8.1Antofagasta earthquake (González and Carrizo, 2003). The cracks herestrike about 020°, parallel to the edge of the salar and to Salar delCarmen fault. More recently, we have observed on the MejillonesPeninsula surface cracks that formed during the 2007 Tocopillaearthquake.

An area of extensive surface cracking occurs to the north–northwest of the Mantos Blanco mine. The cracks occur along, butare much more regionally extensive than, an unnamed north–northeast-striking fault. This fault is shown on Delouis and others'map, butwith the incorrect sense of vertical displacement (their Fig. 2,Delouis et al., 1998); in fact, the southeast side is down-droppedrelative to the northwest side and the dip of the fault is unknown.Surface cracks in this area (Fig.11) strike dominantly NNWbut there isa second set which strikes NE (Loveless et al., 2009). As elsewhere, thecracks are best developed in gypsum indurated saline soils (“gyp-crete”) but they can be found cutting bedrock as well (Fig. 11, inset).


Aside from locally well-developed surface cracks, there arerelatively few young structures in the region between the Mejillonesand Salar Grande segments (Fig. 1). This area has the highest reliefanywhere north of Antofagasta even though the bedrock geology isnot significantly different from segments to the north or south.

4.3. Salar Grande segment

The Salar Grande segment (Fig. 12) has received, perhaps, themostintense recent study of any part of the northern Chile CoastalCordillera north of the Mejillones Peninsula (e.g., González et al.,1997; Reijs and McClay, 1998; Chong et al., 1999; Oncken et al., 1999;González et al., 2003a,b, 2008; Allmendinger et al., 2005; Lovelesset al., 2005; Carrizo et al., 2008). Here, the complete suite of CoastalCordillera structural styles is well represented, including EW reversefaults, NS normal faults and NNW trending strike slip faults, as well asthe most extensive development of surface cracks anywhere in theregion.

4.3.1. Northern terminus of the Atacama fault zone and related structuresThe northernmost on-shore segment of the Atacama fault zone in

this area has a notable fault line scarp on both sides of the Río Loa, yetdisplays little Late Cenozoic offset in the walls of the river canyonitself. The fault does not appear to produce any observable offset inthe EW striking Cerro Aguirre fault system to the south of the river(Fig. 12). To the north, the Salar Grande segment of the Atacama faultzone produces a spectacular scarp up to 30 m high in the pure halitesurface of the salar (Fig. 13); in this area several faults of the EWChuculay system appear to truncate against the structure. To thenorthwest of the salar, the fault produces consistent offsets ofdrainages in a right lateral sense. Offset and abandoned alluvial fanshave surface ages as young as 2.17 Ma (Carrizo et al., 2008). To the



Fig. 6. Map of the Mejillones segment, showing some of the faults and localities described in the text: A–Antofagasta; B–Bandurias fault; CH–Caleta Herradura fault; M–Mejillonesfault; SdC–Salar del Carmen fault; nSdC–northern Salar del Carmen fault; MB–Mantos Blanco. Tick marks are on the downthrown side of faults. Shaded relief and topographiccontours the same as in Fig. 1. White dashed line is the 50 km contour on the depth to the subducted plate. Black dashed line shows location of cross section in Fig. 9. Black dots areaftershocks of the 2007 Tocopilla earthquake.

Fig. 7. Panorama of the Caleta Herradura half graben, looking WSW. The sea cliffs are 60–70 m high.


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Please cite this article as: Allmendinger, R.W., González, G., Neogene to Quaternary tectonics of the coastal Cordillera, northern Chile,Tectonophysics (2009), doi:10.1016/j.tecto.2009.04.019


Fig. 8. (A) Equal area, lower hemisphere projection of minor normal faults in thehanging wall of the Caleta Herradura half graben. Arrows on the great circles show themotion of the hanging wall. The equal numbers of synthetic (east-dipping) andantithetic (west-dipping) faults indicates that a vertical simple shear (“Chevron”)kinematic model is most appropriate. (B) Listric rollover normal fault model fit to thetop of the basement surface, dips of Miocene and Pliocene strata along the beach, andthe dip of the main fault plane at sea level. Solid fault trace is for the vertical simpleshear model; dashed trace is for synthetic shear model.


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west of the Salar Grande fault, the sub-parallel Chomache–BahíaBlanca–Antenna fault system also displays evidence of right lateraldisplacement of modern drainages (González et al., 2003a,b; Carrizoet al., 2008).

4.3.2. EW reverse faultsReverse faults that strike approximately E–W and have dip slip

movement characterize the Coastal Cordillera from the Cerro Aguirrestructure just south of the Río Loa (21.6°S) to the Quebrada deCamarones at the north end of the Atajaña segment (19.2°S). Thesestructures in the Salar Grande region were originally considered to benormal faults (González et al., 1997; Reijs and McClay, 1998), butsubsequent field work, particularly in the regions where the faults cutthe Coastal Escarpment, demonstrated that they are reverse faultswhich dip moderately to the north or south (Allmendinger et al.,2005). Several of the structures (Chuculay, Aguirre, Barranco Alto) cutand offset small, and now inactive paleodrainages; the channels areoffset vertically but not horizontally (Fig. 13), demonstrating that theyhave pure dip-slip motion which is consistent with fault slip data(Allmendinger et al., 2005; Carrizo et al., 2008). In general, the EWreverse faults cut and offset a peneplained surface that has yieldedsurface ages between 10 and 20 Ma (Carrizo et al., 2008) so the faultsare post-early Miocene in age. Allmendinger et al. (2005) used growthstrata at Barranco Alto to show that the structure was growing by5.6 Ma; it probably also offsets the Pleistocene wave-cut platform onthe coast. Carrizo et al. (2008) dated an offset paleochannel in theeastern part of the Chuculay system at about 4 Ma.

Chuculay, the largest EW reverse fault in the region, is notable forextensive development of surface cracks localized in the hanging wallblock. Using numerical modeling and mapping of ~30,000 cracks,González et al. (2008) demonstrate that crack formation is directly


related to the stretching of the uplifted hanging wall block as it foldsabove a propagating tip line or when the fault ramps onto the surface.

4.3.3. North-striking normal and reverse faultsWest of the Salar Grande, young north-striking faults have either

normal or reverse offset. The Punta de Lobos fault (Fig. 12), like mostnormal faults in the region, dips 78° E and has clear normaldisplacement. A 3.9 Ma tuff, embedded in alluvial fan gravels in thehanging wall, but resting on bedrock on the side of an active valley inthe footwall, has been offset by ~1.5 m; thus most of the normaldisplacement predates the ash, but a small amount of normaldisplacement postdates the ash. The same fault produces a scarp,also down to the east, on the cliff face of the Coastal Escarpment(González et al., 2003a,b), indicating that it has remained active untilvery recently. These relations indicate either a very slow continuousslip rate or a very long recurrence interval for slip events on this fault.

In contrast to the Punta de Lobos fault, both the Hombre Muertoand the Geoglifos faults exhibit a mountain bounding scarp morphol-ogy in which the basin is uplifted relative to the mountains on faultsthat dip towards the basin (Carrizo et al., 2008). Likewise, at thewestern border of the Salar Oficina Gloria (Fig. 12), an N–S strikingfault produces the uplift of the basin relative to the mountain front.This fault related morphology suggests that the first order topographywas produced by mountain bounding normal faults, and the mostrecent movements on those faults have been in a reverse sense.

4.3.4. Open tectonic cracksMore than 50,000 open tectonic cracks have been measured on

1 m resolution IKONOS imagery in the Salar Grande areas (Gonzálezet al., 2008; Loveless et al., 2005). Those associated with Chuculay eastof the salar have already been described. The cracks west of the salarhave a strong NNW preferred orientation with the mean directionabout 20° more northerly than the dominant NW strike of faults in theregion. The extensional strain due to cracks is 1–2%, orientedstatistically parallel to the direction of modern plate convergence.Within a few hundred meters of fault scarps, extension due tocracking can increase to 5–10% (Loveless et al., 2005).

4.4. Iquique segment

Within the Iquique segment, the overall trend of the CoastalCordillera changes from an azimuth of ~000° to the south to ~340° tothe north. This change in orientation occurs more-or-less where thetopographic symmetry plane of the Central Andes crosses the CoastalCordillera (Fig. 1). To the north of this axis, the morphologic CoastalCordillera becomes narrower as the structural province trendsoffshore. The narrow, littoral, Pleistocene wave cut platform is almosttotally absent north of the city of Iquique. The north- to northwest-striking structures that control the morphology to the south are verymuch diminished in importance in the Iquique segment and farthernorth as the Atacama fault zone is here located offshore.

The Iquique segment (Fig. 14) does have abundant ~EW-strikingreverse faults, including some of the best exposed of any in the region.Multiple structures with EW fault scarps at the top of the coastalescarpment can be traced down to sea level where they deform thelittoral platform on which the city of Iquique is built. For example, a40 m high EW fault scarp separates the northern part of the city(the “Zofri” zone) from downtown, and the Cavancha Peninsula isformed by another EW structure. But, the most spectacular exposure(C. Marquardt, pers. comm. 2005) occurs at the southern entrance tothe city where the “Los Pacos” reverse fault is well exposed in the 10mhigh sea cliff (Fig. 15). The 30° north-dipping fault places alteredMesozoic volcanic rocks over coarse grained beach deposits that haveyielded late Pleistocene fossils. As the fault plane exits the basementrocks of the footwall, it splays upward into a triangular zone ofdeformation reminiscent of trishear kinematics (Allmendinger, 1998;



Fig. 9. Block diagram and cross-section (at 23°S latitude) looking south of the structures in the Mejillones segment. Topography above sea level is exaggerated by 2×. Section belowsea level is 1:1. Red dots are aftershocks of the November 2007 Tocopilla earthquake. (For interpretation of the references to colour in this figure legend, the reader is referred to theweb version of this article.)

Fig. 10. (A) Topographic bench along the northern Salar del Carmen fault marked by white arrows. (B) Detail of reverse fault relationship showing folding of a Pliocene(?) tuff over abasement block.


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Fig. 11. (A) Image modified from Google Earth showing cracks northwest of Mantos Blanco. (B) Field photo of Jack Loveless in a 2 m wide crack in bedrock.


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Erslev, 1991). The 125,000 year old wave cut surface in the hangingwall is obviously folded and the beach deposits onlap the forelimb.

4.5. Atajaña–Pisagua segment

The Atajaña–Pisagua segment (Fig. 14) has a few significant coast-parallel normal faults, but the primary neotectonic features are a suiteof large, widely spaced southward-dipping reverse faults. The faultsstrike ENE so that they remain perpendicular to the WNW-trendingCoastal Cordillera and die out at the western margin of the CentralValley.

The Atajaña scarp has more than 500 m of vertical relief. At theintersection of the fault and the coastline, a narrow marine platformhas been uplifted and tilted. This platform, like similar structures inthe Iquique segment, attests to youngmotion on the Atajaña fault. Thefault probably originated as a Mesozoic normal fault that hassubsequently been inverted in Late Cenozoic time: the hanging wallis composed of Cretaceous redbeds whereas the footwall containsJurassic volcanic rocks (Anonymous, 2003). Caprio (2007) studied theFalla Blanco structure, a subsidiary scarp to the south of the mainAtajaña scarp. She showed that Falla Blanco and others in the Atajañafamily formed initially as small separate fault segments and only laterin their history have they coalesced into a single structure, thusexplaining the irregular surface traces of all of the faults. She alsoshowed that Falla Blanco, as well as the eastern part of Atajaña andmost of the other smaller scarps in the system, are fold scarps ratherthan scarps of emergent faults.

We first noted the importance of fold scarps in the regionalmorphology of the Coastal Cordillerawhile studying the eastern end ofthe Pisagua scarp to the south of Atajaña (Fig. 14). The fold forms arewell exposed in the Quebrada Tiliviche and other drainages. Allmen-dinger et al. (2005) dated a tuff that was folded over the scarp at3.5 Ma; most of the relief of the scarp has accrued since that time.Pisagua, like Atajaña, also has an isolated uplifted Pleistocene marineterrace where the fault intersects the coast, indicating that movementhas continued into recent times. A crustal earthquake occurred inMarch 2007 at a depth of ~26 km,with an epicenter about 20 km southof the town of Pisagua; the south dipping nodal plane has the samestrike as, and projects to the surface at, the Pisagua fault scarp. The


earthquake was too small (M 5.7) to produce a surface rupture atthat depth, but a month after the event, we observed in the field finesurface cracks with vertical walls in unconsolidated surface materialsthat could be traced for tens of meters and have apertures of a fewmillimeters. These new cracks commonly occurred in the surfacedepression of older cracks with very rounded and degraded walls,suggesting that new cracks are reactivating old crack traces.

5. Discussion: a taxonomy of structural styles

For more than 600 km along strike, the Coastal Cordilleramorphological province tracks the geometry of the subducted plate:the eastern margin of the province is consistently located 60±5 kmabove the subducted Nazca plate (Fig. 1). At Arica, the subducted plateis more than 60 km deep and the Coastal Cordillera is entirelymissing.Onemight argue that, because the inland border of the province is eastof the 50 km depth limit of significant interplate seismic activity, theprovince is not related to coupling of Nazca and South America.However, modeling of the upper plate elastic strain field associatedwith the interseismically locked plate boundary shows significantstrain accumulation extending well east of the zone of seismiccoupling (Fig. 16). Likewise, GPS data demonstrates that significantcoseismic elastic strain release during the 1995 Antofagasta Mw 8.1earthquake occurred as far east as the 75 km depth to subducted platecontour (Klotz et al., 1999). Thus, it is likely that all of the LateCenozoic structures in the region are a direct result of the long termcoupling between the subducted Nazca Plate and the overriding SouthAmerican plate.

The Late Cenozoic structures of the Coastal Cordillera describedabove can be grouped into four distinct types of structures (Fig. 17):(1) margin perpendicular reverse faults that strike approximatelyeast–west; (2) north-striking normal fault zones, some of which showseveral meters of reactivation as reverse faults; (3) reverse faultsstriking NNW to NW; and (4) surface cracks of both tectonic and non-tectonic origins. We have direct evidence from recent earthquakesthat the reverse faults of group 1 form during the interseismic part ofthe plate boundary cycle and that some surface cracks form during thecoseismic part of the cycle. For the rest of the structures, only indirectevidence and numerical modeling point to their origins.



Fig. 13. Shaded relief detail of the Chuculay and Salar Grande fault systems, based on a 20 m digital elevation model by Yu and Isacks (1999).

Fig. 12.Map of the Salar Grande segment, showing some of the faults and localities described in the text: A–Aguirre fault; BA–Barranco Alto fault; Ch–Chuculay fault system; G–Geoglifosfault; ChM–Chomache fault; PL Punta de Lobos fault; SGF–Salar Grande fault; RL–Río Loa. Sawteeth shown on the upper plate of reverse faults Shaded relief and topographic contours thesame as in Fig. 1. White dashed line is the 50 km contour on the depth to the subducted plate. Inset image modified from Google Earth shows an oblique view of the Oficina Gloria salarbounded on the west by a normal fault reactivated as a reverse fault (X), an EW reverse fault (Y) which has offset and uplifted the paleochannels (Z).


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Fig. 14. Map of the Iquique and Atajaña–Pisagua segments, showing some of the faults and localities described in the text: A–Atajaña fault; FB–Falla Blanco; LP–Los Pacos fault; P–Pisaguafault; Z–Zofri fault. Reverse faults shownwith sawteethon theupper plate.Double arrow indicates a fold scarp.Normal faults shownwith tickmarks ondownthrownside. Shaded relief andtopographic contours the same as in Fig. 1. NNW trending, labeled dashed lines are contours on the depth to the subducted plate. Focal mechanism plotted at epicenter of M 5.7 crustalearthquakeat ~26kmdepthwhichoccurredon24March2007. Theheavynodal planehas samestrike as, andprojects to the surface at, thePisagua fault.Oroclinal axis fromGephart (1994).Inset map shows the structures that cut the littoral platform on which the city of Iquique is located. CP–Cavancha Peninsula; AH–Alto Hospicio.


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5.1. Margin-perpendicular reverse faults and margin curvature

The margin perpendicular reverse faults are present only between19°S and 21°40' S latitude and are symmetrically distributed on eitherside of the axis of the Bolivian Orocline (Allmendinger et al., 2005).The faults have pure dip slip motion and produce about 1% marginparallel shortening. They are restricted to the Coastal Cordillera anddie out into fold scarps on the west side of the central valley. Thesestructures have been active for at least the last 6million years and they


are still active today, as shown both by the deformed Pleistocenewavecut terrace in several locations and by the March 2007 crustalearthquake just south of Pisagua (Fig. 14). A very rough estimate of therate of shortening due to these structures is 0.2–0.5 mm/year.

The Pisagua earthquake showed that margin-parallel shorteningcan occur during the interseismic part of the seismic cycle.Furthermore, elastic modeling (Bevis et al., 2001) demonstrates thatinterseismic elastic strain related to a curved plate boundary which isconcave towards the subducted plate should produce a component of



Fig. 15. Photograph of the Los Pacos fault at the southern entrance to Iquique. Note the onlap of the late Pleistocene beach deposits on the forelimb and the folding of the wave-cutterrace in the hanging wall. Thewhite box shows the location of the detailed photograph to the right which shows the fault zone and the growth stratawith clasts of the hanging walllithologies.


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margin-parallel shortening. The ideal elastic model for a fully lockedAndean plate boundary gives−3.3±0.15 nstrain/yearmargin parallelshortening in the Coastal Cordillera (Bevis et al., 2001). Assuming veryroughly that the margin parallel shortening due to the EW reversefaults started sometime between 15 and 3 Ma, the permanent longerterm margin parallel shortening rate is −2±1.3 nstrain/year; thus,the calculated interseismic elastic strain and the long term permanentstrain are very similar. A similar pattern has been demonstrated inCascadia, another concave convergent margin (Johnson et al., 2004;McCaffrey et al., 2007) with little in common with northern Chilebesides the curvature. The primary remaining question is why themargin-parallel elastic strain is not fully recouped during subductionzone slip events but instead is converted into permanent upper platedeformation.

5.2. Margin-parallel normal faults

The margin-parallel normal faults are best developed in the vicinityof Antofagasta and theMejillones Peninsula,wherehorizontal extensionrates vary from 0.025 to 0.18 mm/year, and diminish in importancenorth of Tocopilla. Significant normal faulting does not exist east of the50 km depth contour. The majority of the normal faults dip to the eastand, where present, hanging-wall growth strata suggest that the faults


have been active throughout much of the Neogene. Some normal faultzones in the Antofagasta segment display incipient reactivation asreverse faults (Paposo and northern Salar del Carmen faults) and moreextensive reactivation can be seen for normal faults in the Salar Grandearea. We have insufficient data to determine whether the normal faultsconvert permanently to reverse faultswith timeor if normal and reversefaulting occur repeatedly throughout the life of the structure, with finitehorizontal extension dominating overall. The fact that reactivationappears common in the region that is 35–50 km above the subductedplate would appear to favor the former possibility.

How the normal faults relate to the seismic cycle is not so clear.Margin-perpendicular extension can occur during both the inter-seismic elastic elastic deformation (Fig. 16) and the coseismic parts ofthe cycle (Klotz et al., 1999). In the Antofagasta/Mejillones area, thezone of extension extends from the Salar del Carmen fault on the eastto just east of the trench (von Huene and Ranero, 2003). Only part ofthis transect lies within the regions of interseismic flexural extension,regardless of whether one chooses 38 or 50 km as the extent of down-dip locking for interplate seismic coupling. Thus, it seems likely thatsome normal faults move during major plate boundary thrust eventsand some move during the interseismic part of the cycle.

Some authors have suggested that normal faults are due to longterm subduction erosion (Armijo and Thiele, 1990; von Huene and



Fig. 16.Model of interseismic elastic strain accumulation for a convergent plate boundary which is fully locked to either 38 km (solid line) or 50 km (dashed line). Shaded areas showregions of extension which are due to interseismic elastic flexure. Figure from Loveless (2007).


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Ranero, 2003). If the Nazca and South American plates are mostly orcompletely locked during the interseismic part of seismic cycle, thensubduction erosion could only happen coseismically (Fig. 17). In thiscase, any normal faults that are produced by subduction erosionwould be predominantly coseismic as well.

Fig. 17. Taxonomy of forearc structural styles and their relatio


5.3. NW to NNW strike–slip faults

Significant strike–slip appears to be limited to the relatively smallnumber of west–northwest to north–northwest striking faultsvirtually all of which have right lateral displacement. The kinematics

ns to the interplate seismic cycle. See text for discussion.



Fig. 18. Image modified from Google Earth showing tectonic and non-tectonic cracks. The latter form the fine polygons, whereas the former can be traced for hundreds of meters inlength and have a rough preferred orientation. The heavy dark lines are drainages; note that the tectonic cracks cross the drainages.


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of the WNW-striking Salar de Navidad and Barazarte faults southeastof Antofagasta are inconsistent with interseismic shortening acrossthe margin but consistent with coseismic elastic rebound as well asmargin parallel shortening. Nonetheless, we have no direct evidence ofwhen they lastmoved.Northof theRío Loa, theSalarGrande, Chomache,and Antena faults strike NNW. They are nearly perpendicular to theconvergence direction but kinematically consistent with interseismicshortening parallel to the margin (Carrizo et al., 2008).

5.4. Surface cracks

In a hyperarid environment, surface cracks can formdue to a varietyof mechanisms, both non-tectonic and tectonic (Fig. 18). The formerare those due largely to surface processes such as saline soil formationand mass wasting. Pedogenic cracks form due to salt wedging andexpansion and contraction of gypsum indurated crusts (Tucker, 1978).Though ubiquitous in the field, cracks produced by this process aresufficiently small that they seldom can be resolved on the 1 m–2.5 mresolution imagery that we currently use to map the cracks. Oneexception to this rule is the region south of the Río Loa where one canresolve both tectonic and non-tectonic cracks in the imagery (Fig. 18);the non-tectonic cracks tend to form equidimensional polygonswhereas the tectonic cracks are hundreds of meters long, have apreferred orientation, and cross paleovalleys. Non-tectonic cracks mayalso formdue to slidingonunstable slopes, especially near canyons andsteep hillsides. Soil cracks probably form slowly over long periods oftime and thus form during interseismic and coseismic parts of thecycle. If is likely that cracks due to slope instabilities formprimarily dueto shaking during earthquakes (e.g., Marquardt et al., 2006).

All types of cracks, including the tectonic cracks described below,can be enhanced by surface processes (Keefer and Moseley, 2004).Water runoff during the infrequent precipitation events that affect theregion dissolve the thick salts filling the cracks, leaving meter scalesubsurface voids, and leading to roof collapse. Once a crack is formed,regardless of mechanism, soil salts fill the crack and contribute tofurther wedging.

Most of the N55,000 surface cracks that we have mapped onIKONOS and Google Earth imagery (Carrizo et al., 2008; Gonzálezet al., 2008; Loveless et al., 2005, 2009) are primarily tectonic; that is,


they reflect deep seated, crustal scale processes rather than surfaceeffects. Commonly, tectonic cracks are filled with gypsum, windblownsand, or both. The gypsum is usually vertically banded, with theopening occurring in the center of the structure. Thus, we interpretthat most cracks have a protracted history of opening and filling.Where exposures permit, the cracks can be seen to penetrate bedrockfor at least 9–12 m and are probably considerably deeper.

We distinguish three types of tectonic cracks: (1) those that formduringupper crustal faulting that is unrelated to the seismic cycle on theplate boundary. The cracks associated with the Chuculay reverse faultsystem are an example (González et al., 2008). (2) Cracks that formduring plate boundary earthquakes and are due to focusing of surfacewaves by topographic scarps, basinmargins, and zones of weakness likefaults. Numerical modeling by Loveless (2007) has shown that thefocusing is only important within a few hundred meters of the fault orscarp. Cracks thatopenedon thewestern side of Salardel Carmenduringthe 1995 Antofagasta earthquake were probably due to surface wavefocusing due to the edge of the salar basin and the adjacent Salar delCarmen fault. Cracks that lie at greater distances from known structuresand scarps are considered to be (3) features that form coseismically,probably due to extension during elastic rebound of the upper plate.Modeling by Loveless (2007) of dynamic wave propagation during anAntofagasta-type event in a uniform half space shows that the dynamicprincipal horizontal extension axes and the static principal extensionaxes have the same orientation.

Ample evidence confirms that coseismic processes are importantin crack formation. The most direct evidence is the observed crackingthat occurred during the 1995 Antofagasta, 2001 Arequipa (Keefer andMoseley, 2004), 2005 Tarapacá, and 2007 Tocopilla earthquakes. Eventhough the latter two were smaller than the two former ones, newfine cracks were observed in the center of old cracks after both events(Fig. 19). Because the cracks are long lived, and because of theextraordinary preservation of these delicate structures in thehyperarid Atacama Desert, their distribution may provide insightinto seismic segmentation of the plate boundary.

As documented by Loveless et al. (2009), the cracks displayregionally consistent orientation locally and vary systematically inorientation regionally. At the northern end of the area described here,the cracks from groups 2 and 3, above, are bimodal, striking NW and



Fig. 19. Reactivation of an older surface crack during the 2007 Tocopilla earthquake. The new crack is in the unconsolidated coarse sandy fill of the older crack. The older crack wallshave been partially highlighted with dashed white lines.


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NE. In the middle of the area, they tend to be unimodal and strikeNNW, and at the south end of the area, they are, again, bimodal withboth NE and NW strikes well represented. Loveless et al. (2009) haveproposed that the coseismic cracks can be used to define long-livedsegments. Their inversion of crack orientations for a best fittingrupture zone for the Iquique segment indicates thatmost of the slip onthe plate boundary should occur north of the Río Loa. There is a broadzone of transition between Río Loa and Mejillones to the south wherethere is high topography and relatively few cracks and those that existdisplay bimodal orientations. The large, but not great, November 2007Tocopilla earthquake occurred in this gap, hinting at the intriguingpossibility that this segment may not experience numerous greatearthquakes. Likewise there is a broad zone of transition at the northend of this segment in the Arica area, which also displays bimodalorientations.

6. Conclusions

The Late Cenozoic tectonics of the Coastal Cordillera of northernChile reflect processes related to the seismic coupling between thesubducted Nazca Plate and the overriding South American Plate.Although these processes probably occur in all eroding convergentmargins around the globe, only in northern Chile is the recordpreserved due to the hyperarid climate of the region. Our study of theCoastal Cordillera has yielded the following conclusions:

1. Margin-parallel shortening on ~EW reverse faults is probablyrelated to the curvature of the margin, which is concave towardsthe subducted plate (Bevis et al., 2001). Additional shorteningcould also occur due to oblique subduction (McCaffrey, 1996).

2. Margin-perpendicular extension probably occurs during bothinterseismic and coseismic parts of the plate boundary seismiccycle. The interseismic extension is due to flexure of the margin,and coseismic extension is due to elastic rebound and subductionerosion.

3. Surface cracks form due to a number of mechanisms, but aparticularly important one is coseismic extension. Regional varia-tions in orientations of the cracks suggest that rupture segments ofthe plate boundary persist over long periods of time, perhapsmillions of years (Loveless et al., 2009).


4. Upper plate fault zones are probably very weak and prone to slip indifferent directions due to minor changes in the stress field.

5. Forearc deformation rates are very slow, with average fault sliprates less than 0.5 mm/year and strain rates on the order of 1–5 nstrain/year.

A great deal more will be learned about these processes when theCoastal Cordillera is completely covered by a dense network ofcontinuous GPS stations. This will allow us to understand much morefully the role of interseismic deformation, particularly interseismicflexure. Unfortunately, we will also learn much more about theseprocesses once the next great Iquique earthquake occurs. It has beenmore than 130 years since the last great earthquake and recent earth-quakes to the north (2005 Arequipa) and south (1995 Antofagasta,2007Tocopilla) of the Iquique segment suggest thatwewill not have towait long.

Acknowledgments

We are indebted to numerous colleagues and students who, overthe last decade, have shared their knowledge of the Coastal Cordillera,as well as, more often then not, their good humor in the field. Inparticular, we wish to thank our students Jack Loveless and DanielCarrizo whose work we have drawn on heavily in this paper.Additionally, we are grateful to José Cembrano, Greg Hoke, HollyCaprio, Terry Jordan, Matt Pritchard, Bryan Isacks, Tibor Dunai,Alejandro Macci, Francisco Gomez, Joaquin Cortés and Jacob Espina.We are grateful to Pia Victor and Jonas Kley for their careful reviews ofan earlier version of this manuscript. Allmendinger and his students'work in northern Chile has been funded by the U. S. National ScienceFoundation through grants EAR 0087431, EAR-0337496, and EAR-0738507. González and his students' work was funded by theFONDECYT projects 1040389 and 1085117.

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